Biogeosciences, 12, 6881–6896, 2015

www.biogeosciences.net/12/6881/2015/

doi:10.5194/bg-12-6881-2015

© Author(s) 2015. CC Attribution 3.0 License.

Carbonate saturation state of surface waters in the Ross Sea and

Southern Ocean: controls and implications for the onset of

aragonite undersaturation

H. B. DeJong, R. B. Dunbar, D. Mucciarone, and D. A. Koweek

Department of Earth System Science, Stanford University, Stanford, CA, USA

Correspondence to: H. B. DeJong (hdejong@stanford.edu)

Received: 6 May 2015 – Published in Biogeosciences Discuss.: 8 June 2015

Revised: 30 October 2015 – Accepted: 23 November 2015 – Published: 2 December 2015

Abstract. Predicting when surface waters of the Ross Sea

and Southern Ocean will become undersaturated with re-

spect to biogenic carbonate minerals is challenging in part

due to the lack of baseline high-resolution carbon system

data. Here we present ∼ 1700 surface total alkalinity mea-

surements from the Ross Sea and along a transect between

the Ross Sea and southern Chile from the austral autumn

(February–March 2013). We calculate the saturation state of

aragonite (�Ar) and calcite (�Ca) using measured total al-

kalinity and pCO2. In the Ross Sea and south of the Polar

Front, variability in carbonate saturation state (�) is mainly

driven by algal photosynthesis. Freshwater dilution and cal-

cification have minimal influence on � variability. We esti-

mate an early spring surface water �Ar value of ∼ 1.2 for

the Ross Sea using a total alkalinity–salinity relationship and

historical pCO2 measurements. Our results suggest that the

Ross Sea is not likely to become undersaturated with respect

to aragonite until the year 2070.

1 Introduction

Atmospheric CO2 concentrations have increased by 40 %

since preindustrial times to∼ 400 ppm today and could reach

936 ppm by the year 2100 (IPCC AR5 WG1, 2013). Due to

oceanic uptake of CO2, surface ocean pH is already 0.1 units

lower than preindustrial values and is projected to decrease

by another 0.3–0.4 units by the end of the century, equivalent

to a 50 % decrease in carbonate ion (CO2−3 ) concentrations

(Orr et al., 2005). Even after CO2 emissions are halted, it will

take thousands of years before the surface ocean pH returns

to preindustrial levels (Caldeira and Wickett, 2003; Archer et

al., 2009).

The saturation state (�) of seawater with respect to a spe-

cific calcium carbonate (CaCO3) mineral (aragonite, calcite,

or magnesium calcite) is defined as

�=

[Ca2+][CO2−

3

]Ksp

, (1)

where Ksp is the solubility product constant for the specific

CaCO3 mineral and depends on salinity, temperature, and

pressure (Mucci, 1983). Aragonite is ∼ 1.6 times more sol-

uble than calcite at 0 ◦C whereas the solubility of magne-

sium calcite varies depending on the mole fraction of mag-

nesium ions (Dickson, 2010). �Ar represents the saturation

state of aragonite and �Ca represents the saturation state of

calcite. � < 1 represents undersaturation where dissolution is

thermodynamically favorable and � > 1 represents supersat-

uration where precipitation is favorable. Most surface wa-

ters of the global oceans are currently supersaturated with

respect to CaCO3 (Feely et al., 2009). However, for some

species including coccolithophorids, foraminifera, and tropi-

cal corals, decreasing CO2−3 concentrations can decrease cal-

cification rates even in supersaturated conditions (Riebesell

et al., 2000; Moy et al, 2009; Andersson et al., 2011).

The Southern Ocean is especially vulnerable to ocean

acidification (OA) due to its relatively low total alkalin-

ity (TA) and because of increased CO2 solubility in cold wa-

ter. In addition, Antarctic continental shelves have insignifi-

cant sedimentary CaCO3 to buffer against OA (Hauck et al.,

2013). Modeling studies predict that surface waters in the

Southern Ocean may start to become undersaturated with

Published by Copernicus Publications on behalf of the European Geosciences Union.

6882 H. B. DeJong et al.: Carbonate saturation state in the Ross Sea and Southern Ocean

respect to aragonite by 2050 and be fully undersaturated

by 2100 (Orr et al., 2005; Feely et al., 2009). McNeil and

Matear (2008) have suggested that wintertime aragonite un-

dersaturation in the Southern Ocean may begin as early as

2030.

OA-induced decreases in � have potentially serious con-

sequences for Antarctic food webs. In the Ross Sea the

aragonitic shelled pteropod Limacina helicina is a domi-

nant zooplankton that can reach densities of 300 individu-

als m−3 (Hopkins, 1987; Seibel and Dierssen, 2003; Hunt et

al., 2008). Pteropods are important prey for notothenioid fish,

which in turn are major prey for penguins, seals, and whales

(Foster and Montgomery, 1992; La Mesa et al., 2000, 2004).

Pteropods may also be important contributors to the biologi-

cal pump (Collier et al., 2000; Accornero et al, 2003; Manno

et al., 2010). Orr et al. (2005) found that the shell of a sub-

arctic pteropod started to dissolve within 48 h when placed

in waters with the level of aragonite saturation expected to

occur in the Southern Ocean by 2100. Severe dissolution pit-

ting was observed on live pteropods that were collected from

the upper 200 meters in the Atlantic sector of the Southern

Ocean, from waters that were near undersaturation with re-

spect to aragonite (Bednaršek et al., 2012).

Other organisms in the Southern Ocean may be nega-

tively impacted by OA include krill (Kawaguchi et al., 2013),

foraminifera (Moy et al., 2009), sea urchins (Sewell and Hof-

mann, 2011), deep sea hydrocorals (Shadwick et al., 2014),

sea stars (Gonzalez-Bernat et al., 2013), bivalves (Cummings

et al., 2011), and brittle stars (McClintock et al., 2011).

Conversely non-calcareous phytoplankton may benefit in the

Ross Sea in a high pCO2 world, especially the larger diatom

Chaetoceros lineola (Tortell et al., 2008; Feng et al., 2010).

There are only a few surface carbon system data sets from

the Ross Sea (Bates et al, 1998; Sweeney et al., 2000b; San-

drini et al., 2007; Long et al., 2011; Mattsdotter Björk et al.,

2014; Rivaro et al., 2014; Kapsenberg et al., 2015) that can be

used to establish baselines in order to understand the relative

importance of physical, chemical, and biological processes

that drive the large spatial and seasonal variability of �. With

no winter � measurements, it is challenging to predict when

the Ross Sea will become undersaturated with respect to

aragonite and calcite. A model by McNeil et al. (2010) sug-

gests that winter surface waters in the Ross Sea will become

undersaturated with respect to aragonite by the year 2045

since sea ice, upwelling of deep water, and short residence

times prevent these surface waters from reaching equilibrium

with the atmosphere. However, McNeil et al. (2010) indi-

rectly estimated surface winter �Ar values by using limited

carbon system data from the spring (Sweeney et al., 2000b).

We present ∼ 1700 underway TA measurements from the

surface waters of the Ross Sea and along a transect across

the Southern Ocean from the Ross Sea to southern Chile.

By combining the underway TA measurements with pCO2

data we characterize the complete carbon system and de-

scribe patterns and controls on � variability. Finally, af-

ter establishing a relationship between salinity and TA, we

use the Lamont Doherty Earth Observatory (LDEO) pCO2

database (Takahashi et al., 2009) (available at http://www.

ldeo.columbia.edu/res/pi/CO2) to provide an independent es-

timate of Ross Sea surface water �Ar in early spring.

2 Study site

The Antarctic Circumpolar Current (ACC) flows from east to

west around the entire Antarctic continent and is composed

of multiple fronts that separate distinct water masses (Rin-

toul et al., 2001). There are three primary fronts – the south-

ern ACC front (SACCF), the Antarctic Polar Front (PF), and

the Subantarctic Front (SAF) (Orsi et al., 1995). Sokolov and

Rintoul (2009) found that these primary fronts are composed

of multiple jets that they label south (S), middle (M), and

north (N). Convergent Ekman transport north of the westerly

wind stress maximum (near the axis of the ACC) downwells

surface water into the ocean interior. Circumpolar Deep Wa-

ter (CDW) upwells south of the wind stress maximum where

it becomes modified into Antarctic surface water (AASW)

(Rintoul et al., 2001).

The cyclonic Ross Sea gyre is located south of the ACC

(Smith et al., 2012). The southern portion of this gyre flows

west along the Ross Sea continental slope and generates in-

trusions of CDW onto the Ross Shelf through the major

troughs (Orsi et al., 2009; Dinniman et al., 2011; Kohut et al.,

2013). In addition, AASW enters the Ross Sea in the east and

flows westward along the Ross Ice Shelf (Orsi et al., 2009).

The Ross Sea is considered a biological hotspot supporting

over 400 benthic species (Smith et al., 2012). During the win-

ter the Ross Sea is mostly covered by sea ice, which begins

to clear in November to form the largest polynya in Antarc-

tica. There are two main phytoplankton blooms in the Ross

Sea. The first bloom begins in late November in the Ross Sea

polynya (Fig. 1a) and peaks in mid- to late December (Ar-

rigo et al., 1999; Arrigo and van Dijken, 2004). In early Jan-

uary, sea ice melts in the western Ross Sea, lowering surface

salinity and increasing stratification (Fig. 1b). As a result, a

secondary diatom bloom forms in the west with productivity

peaking in late January to early February (Arrigo et al., 1999;

Arrigo and van Dijken, 2004) (Fig. 1c).

The Ross Sea phytoplankton blooms account for up to half

of all primary production over the Antarctic continental shelf

(Arrigo and McClain, 1994; Smith and Gordon, 1997; Ar-

rigo and van Dijken, 2003). Photosynthesis reduces the con-

centration of nutrients and dissolved inorganic carbon (DIC)

in the mixed layer, causing � to increase in surface waters

(McNeil et al., 2010). Once the sea ice reforms during au-

tumn and winter, remineralization of organic matter and deep

convective mixing produces a relatively homogeneous water

column, causing surface DIC concentrations to increase and

� to decrease (Gordon et al., 2000; Sweeney et al., 2000b;

Petty et al., 2014).

Biogeosciences, 12, 6881–6896, 2015 www.biogeosciences.net/12/6881/2015/

H. B. DeJong et al.: Carbonate saturation state in the Ross Sea and Southern Ocean 6883

Figure 1. Maps of (a) 6.25 km gridded sea-ice concentration on 1

December 2012 from the University of Bremen, http://www.iup.

uni-bremen.de:8084/amsr2/ (Spreen et al., 2008); (b) sea surface

salinity from NBP 13-02; (c) satellite chlorophyll concentration

on February 2013 from the 9 km Level 3 Aqua MODIS product,

http://oceancolor.gsfc.nasa.gov/cgi/l3; (d) sTA from NBP 13-02;

(e) pCO2 from NBP 13-02; (f) aragonite saturation state (�Ar)

from NBP 13-02.

3 Methods

3.1 Analytical methods

As part of the TRacing the fate of Algal Carbon Export in

the Ross Sea (TRACERS) program, we undertook continu-

ous measurements of surface water TA in the western Ross

Sea aboard the Nathaniel B. Palmer (NBP13-02) from 13

February through 9 March 2013. In addition, from 19 March

to 2 April 2013, we made continuous measurements of sur-

face water TA in transit between the Ross Sea and southern

Chile along the cruise track shown in Fig. 2. Underway TA

measurements were conducted using the shipboard uncon-

taminated continuous flow system with an intake located at

∼ 5 m depth. Seawater from the ship’s underway system was

redirected to the bottom of a 250 mL free surface interface

cup flowing at 2 L min−1 and was drawn from the bottom of

the cup for TA analysis without filtration. The entire system

165o E 180oW 165 o

W

78o S

76o S

74o S

72o S

70o S

(a) (b)

Figure 2. Cruise track (black line) from NBP 13-02. Stations used

in this study (red circles) from (a) TRACERS (NBP 13-02) and

(b) CLIVAR (NBP 11-02). Blue line is the 1000 m isobath.

was automated and relatively unattended. The sampling cy-

cle was every 24 min on a custom-configured Metrohm 905

Titrando equipped with three Metrohm 800 Dosino syringe

pumps (two 50 mL units for sample handling and rinsing and

one 5 mL unit for acid titration). Temperature was measured

at the cup and in the titration cell. We used certified 0.1 N HCl

provided by A. Dickson (Scripps Institution of Oceanogra-

phy) for the potentiometric titrations and TA calculations fol-

low Dickson et al. (2003). Since we consumed the certified

HCl after ∼ 1000 measurements in the Ross Sea, we have

no TA data from the eastern Ross Sea. For the transect to

southern Chile, we mixed our own 0.1 N HCl solution (from

12.1 N HCl, laboratory grade NaCl, and deionized water). We

calibrated TA measurements using certified reference mate-

rials (CRMs) Batch 122 provided by A. Dickson (Scripps

Institution of Oceanography). Our estimated precision for

the underway TA measurements from 68 CRM analyses is

±3 µmol kg−1 (±1 SD).

Outlier TA analyses were identified by taking a running

mean and standard deviation of nine consecutive measure-

ments. A measurement was rejected if (1) the difference be-

tween the measurement and mean was greater than twice the

standard deviation and (2) the difference between the mea-

surement and mean was greater than 6 µmol kg−1. A total of

65 measurements (out of 1716) were rejected.

We collected seawater samples for particulate organic car-

bon (POC) every 2 h from the ship’s continuous flow system

between the Ross Sea and Chile. Following the protocols of

Knap et al. (1996), we filtered 1 to 3 L of seawater through

precombusted Whatman GFC filters and immediately rinsed

these filters with 10 mL of 0.01 N HCl to remove carbonate.

We air-dried the filters before sending them to Stanford Uni-

versity where they were analyzed on a Carlo Erba NA1500

Series 2 elemental analyzer.

Surface pCO2 measurements were made every 3 min us-

ing the LDEO air–sea equilibrator permanently installed on

the NBP (data available at http://www.ldeo.columbia.edu/

res/pi/CO2). The estimated precision is ±1.5 µatm.

www.biogeosciences.net/12/6881/2015/ Biogeosciences, 12, 6881–6896, 2015

6884 H. B. DeJong et al.: Carbonate saturation state in the Ross Sea and Southern Ocean

Underway salinity and sea surface temperature (SST)

were measured continuously by the ship’s thermosalino-

graph (TSG) (Sea-Bird model SBE-45). These variables

were binned into 1 min intervals.

We collected discrete water samples at 85 stations in the

Ross Sea from 13 February through 18 March 2013 (Fig. 1a).

We used a rosette sampler fitted with 24 Niskin bottles and

a Sea-Bird model SBE-911+ conductivity, temperature, and

depth sensor. We also measured salinity on discrete under-

way and hydrocast samples at 25 ◦C using a Guildline 8400

Autosal four-electrode salinometer. The difference between

the Autosal measurements and salinity from the conductivity

sensor was less than 0.02. In this paper we use the hydro-

cast samples to evaluate the controls of seasonal surface �Ar

variability. The water column data will be further analyzed

in upcoming papers.

We collected hydrocast samples for TA and DIC following

the protocols of Dickson et al. (2007) and immediately added

saturated mercuric chloride (<0.1 % by volume). For TA, we

ran each sample within 12 h of collection using a second po-

tentiometric titrator, a Metrohm 855 Robotic Titrosampler

equipped with two 800 Metrohm Dosino syringe pumps (one

50 mL unit for rinsing and sample handling and one 5 mL

unit for acid titration). The samples were prefiltered through

0.45 µm polyvinylidene fluoride filters and the estimated pre-

cision based on the CRMs (n= 108) is ±1.5 µmol kg−1.

We measured DIC on hydrocast samples within ∼ 4 h of

collection without filtration. We acidified 1.25 mL of the

sample using a custom-built injection system coupled to an

infrared gas analyzer (LI-COR LI7000). As described by

Long et al. (2011), the infrared absorption signal versus time

is integrated for each stripped gas sample to yield a to-

tal mass of CO2. Samples were run in triplicate or greater

and were calibrated using CRMs between every 3–4 un-

knowns. Micro-bubbles regularly appeared within injected

samples due to sample warming between acquisition and

DIC analysis. Each integration curve was visually inspected

and integration curves that exhibited evidence for bubbles

were rejected. The estimated precision based upon unknowns

(> 3500 runs) and CRM replicates (n= 855) for cruise NPB-

1302 is ±3 µmol kg−1.

3.2 Carbon system calculations and crosschecks

We calculate � and DIC (hereafter called DICcalc) for un-

derway samples with CO2SYS for MATLAB (Lewis and

Wallace, 1998; van Heuven et al., 2011) with TA, pCO2,

SST, and salinity as input variables. Calculations are only

conducted for pCO2 measured within 3 min of the TA mea-

surement (n= 1034), the average cycle time for the auto-

mated pCO2 measurements. We use the equilibrium con-

stants of Mehrback et al. (1973) as refit by Dickson and

Millero (1987) since previous studies have found that they

are the optimal choice, including for Antarctic waters (Lee et

al., 2000; Millero et al., 2002; McNeil et al., 2007). For the

hydrocast data, we calculate � using TA, DIC, temperature,

and salinity as input variables.

As a means of internal quality control, we use the initial

pH reading from the TA titration as a third carbon system

parameter to crosscheck the accuracy of our �Ar estimates.

�Ar calculated using TA and pCO2 is 0.02± 0.07 greater

than �Ar calculated using TA and pH. In addition, DICcalc

using TA and pCO2 is 2± 7 µmol kg−1 lower than DICcalc

using TA and pH. Finally, measured pCO2 is 4± 14 µatm

lower than pCO2 calculated from TA and pH. These strong

consistencies suggest that our pCO2 and TA measurements

are accurate. Our surface TA and DICcalc measurements ver-

sus latitude for the Southern Ocean are within the ranges of

other studies (Metzl et al., 2006; McNeil et al., 2007; Matts-

dotter Björk et al., 2014).

We compare the TA measurements from the surface hydro-

casts (< 5 m deep) to the underway TA measurements made

while the ship was still on station within ∼ 15 min of when

the surface samples were collected. The underway values are

3± 5 µmol kg−1 higher than the hydrocast TA values.

3.3 Ross Sea and Southern Ocean calculations

The �Ar of surface waters in the Ross Sea increases dur-

ing the austral summer months (McNeil et al., 2010). We

use DIC, TA, SST, and salinity to determine the con-

trols on the seasonal cycle of surface water �Ar. We

normalize DIC and TA to a salinity of 34.5, the aver-

age salinity of the Ross Sea (hereafter called sDIC and

sTA). Due to the deep convective mixing during the win-

ter, we use the average sDIC and sTA concentrations of

hydrocast samples collected from 200 to 400 m to de-

termine winter water values (sDIC= 2221± 5 µmol kg−1,

sTA= 2338± 3 µmol kg−1). While sDIC and sTA concen-

trations below 200 m are influenced by carbon export par-

ticularly in the summer and early autumn, observations show

that sDIC and sTA concentrations are relatively uniform be-

low 200 m across space and a given season (Table 1).

Following Hauri et al. (2013), the change in �Ar of surface

hydrocast samples (upper 10 m) from winter conditions can

be expressed as

1�Ar =∂�

∂DIC1sDIC+

∂�

∂TA1sTA

+∂�

∂T1T +1S�+ Residuals, (2)

where

1S� =∂�

∂S1S+

∂�

∂DIC1DICs

+∂�

∂TA1TAs . (3)

1sDIC and 1sTA are the difference in sDIC and sTA for

each sample from the winter value. The term 1T is calcu-

lated using a winter SST of −1.89 ◦C (per Sweeney, 2003).

1S� represents the total contribution of salinity changes to

1�Ar.

Biogeosciences, 12, 6881–6896, 2015 www.biogeosciences.net/12/6881/2015/

H. B. DeJong et al.: Carbonate saturation state in the Ross Sea and Southern Ocean 6885

Table 1. Mean values for sDIC concentrations below 200 m.

Data source Early spring Spring Summer Autumn

Sweeney et al. (2000) 2226± 3 2233± 3 2237± 3 2233± 5

Long et al. (2011 ) 2224± 5 2225± 4

This paper 2220± 5

Since salinity between 200 and 400 m is variable across

the Ross Sea (Orsi and Wiederwohl, 2009), 1S is calculated

as the difference between the salinity of a surface sample and

the average salinity for samples from that station that are be-

tween 200 and 400 m.

1DICs and 1TAs represent changes to DIC and TA due

to dilution/concentration from freshwater input and sea-ice

processes:

1DICs=

[DIC200−400 ·

(Salinitysurface sample/Salinity200−400

)]−DIC200−400, (4)

1TAs=

[TA200−400 ·

(Salinitysurface sample/Salinity200−400

)]−TA200−400. (5)

DIC200−400, TA200−400, and Salinity200−400 are the average

values for samples collected from 200 to 400 m calculated at

each station.

The partial derivatives quantify the change in �Ar per unit

change in DIC, TA, temperature, and salinity. To determine

the partial derivatives, we calculate �Ar for all hydrocast

samples within the upper 10 m using DIC, TA, temperature,

and salinity as input parameters. We recalculate �Ar after in-

dependently increasing DIC, TA, temperature, and salinity

by one unit. The partial derivatives are the average difference

between the initial �Ar and the recalculated �Ar.

We use the same equations to evaluate the relative impor-

tance of DIC, TA, temperature, and salinity on the variabil-

ity of �Ar from 75 to 55◦ S. For the 1 terms, we calculate

the change in sDIC, sTA, temperature, and salinity from the

mean of the first six underway measurements at 75◦ S. For

Eqs. (4) and (5), instead of using DIC, TA, and salinity values

from 200 to 400 m, we use the mean of the first six underway

measurements at 75◦ S.

4 Results and discussion

4.1 � in the Ross Sea

Underway TA values range from 2268 to 2346 µmol kg−1

(mean= 2314± 16 µmol kg−1). Since TA strongly covaries

with salinity (R2= 0.86, residual ±6 µmol kg−1), the lowest

TA values are located in the west where the salinity is lowest

(Fig. 1b). Values of sTA range from 2336 to 2386 µmol kg−1

(mean= 2360± 7 µmol kg−1) and are influenced by calcifi-

cation/dissolution as well as phytoplankton photosynthesis

since one unit of nitrate drawdown increases TA by one unit

(Brewer and Goldman, 1978) (Fig. 1d).

−0.2

0

0.2

0.4

0.6

0.8

Con

trib

utio

n to

ΔΩ

Ar

Total sDIC sTA Temp Sal pTA

0.572 ± 0.192

0.525 ± 0.180

0.108 ± 0.063

0.001 ± 0.002

−0.012 ± 0.005

−0.002 ± 0.046

Decomposition of ΔΩAr

from winter

Figure 3. Contributions of sDIC, sTA, temperature, salinity, and

PALK to changes in the aragonite saturation state (�Ar) of surface

waters from the winter to early autumn. Error bars represent±1 SD.

Surface pCO2 values range from 162 to 354 µatm and are

lower in the west due to late-season phytoplankton photo-

synthesis (Fig. 1e). Surface �Ar ranges from 1.40 to 2.42

and �Ca ranges from 2.24 to 3.89 (Fig. 2f). The highest �Ar

values are also located in the west. Phytoplankton photosyn-

thesis increases � by both decreasing DIC and increasing

TA.

Spatial and temporal variations in surface water �Ar are

mainly controlled by sDIC in the Ross Sea (Eq. (2), Fig. 3).

The concentration of sDIC decreased by 58± 20 µmol kg−1

from a winter value, causing �Ar to increase by 0.5± 0.2.

In addition, sTA increased by 11± 7 µmol kg−1 during the

preceding summer months, causing �Ar to increase by

0.1± 0.1. Although there was a significant reduction in salin-

ity compared to winter values (0.7± 0.3), �Ar only de-

creased by ∼ 0.01 due to this freshening since both DIC and

TA concentrations were reduced. Lastly, the effect of temper-

ature on �Ar was negligible since the Ross Sea only experi-

ences a 2 ◦C seasonal change in SSTs (Sweeney, 2003).

Two processes can reduce sDIC, calcification, and phy-

toplankton photosynthesis. To evaluate the importance of

calcification, we use time-dependent changes in potential

alkalinity (PALK= sNitrate+ sTA) from a winter value

(2367± 3 µmol kg−1, defined as average value for all sam-

ples between 200 and 400 m). While TA will increase during

photosynthesis due to nitrate drawdown, PALK will be con-

served. Therefore, changes in PALK can be attributed to cal-

cification and dissolution. The average 1PALK from a win-

ter concentration is negligible (0± 5 µmol kg−1); therefore,

calcification appears to be insignificant and the increase in

sTA from winter conditions is largely driven by nitrate draw-

down during photosynthesis. Earlier studies found that calci-

fication contributed to only ∼ 5 % of the total seasonal DIC

drawdown (Bates et al., 1998; Sweeney et al., 2000a). There-

fore, we argue that photosynthesis exerts the dominant con-

trol on sDIC, sTA, and �Ar. While the highest �Ar value

that we observed was 2.4, values up to ∼ 4 have been ob-

served during December–January (McNeil et al., 2010). By

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6886 H. B. DeJong et al.: Carbonate saturation state in the Ross Sea and Southern Ocean

1.2

1.4

1.6

1.8

2

ΩA

r

SACCF−NPF−S

PF−M PF−NSAF−S

SAF−MSAF−N(a)

−2

0

2

4

6

8

Tem

p [ ° C

]

(b)

320

340

360

380

400(c)

pCO

2 [μat

m]

−75 −70 −65 −60 −550

50

100

150

200(d)

PO

C [μ

g C

L−

1 ]

Latitude [ °S]

Figure 4. Surface water properties from a Southern Ocean tran-

sect, 20 March–2 April 2013: (a) aragonite saturation state (�Ar),

(b) SST, (c) pCO2, and (d) particulate organic carbon. The loca-

tions of the Subantarctic Front (SAF), the Polar Front (PF), and

the southern Antarctic Circumpolar Current Front (SACCF) from

Sokolov and Rintoul (2009) are indicated (gray lines).

the time we arrived in the Ross Sea, surface sDIC concen-

trations would have already increased relative to the summer

due to enhanced air–sea CO2 fluxes (Arrigo and van Dijken,

2007), deepening of the mixed layer (Sweeney, 2003), and

remineralization of organic carbon (Sweeney et al., 2000b).

Mattsdotter Björk et al. (2014) also argue that phytoplank-

ton photosynthesis is the major control on surface water �Ar

variability between the Ross Sea and the Antarctic Peninsula

based upon the covariance of �Ar and chlorophyll a. The

largest contributor to seasonal �Ar change in the Chukchi

Sea in the Arctic is also phytoplankton photosynthesis (Bates

et al., 2013). However, unlike the Ross Sea, numerous stud-

ies have also demonstrated aragonite undersaturation of sur-

face waters in parts of the Arctic due to sea-ice melt and river

runoff (Chierici and Fransson, 2009; Yamamoto-Kawai et al.,

2009; Robbins et al., 2013).

4.2 � in the Southern Ocean

The spatial changes in �Ar, SST, pCO2, and POC between

75 and 55◦ S are shown in Fig. 4. We also include the mean

location of the fronts from Sokolov and Rintoul (2009) as

they intersect our cruise track. The lowest �Ar value is 1.25

(�Ca = 2.00) at 75◦ S, corresponding with the highest pCO2

of ∼ 396 µatm. �Ar increases along the transect to reach a

maximum of 1.93 (�Ca= 3.04) at 55◦ S. The changes in �Ar

are not always monotonic. In two regions changes in �Ar

can be attributed to enhanced primary production. Between

74 and 73◦ S, �Ar first increases and then decreases by∼ 0.1.

This corresponds with a 40 µatm drop and then rise in pCO2.

−0.4

0

0.4

0.8

1.2

(a)

Con

trib

utio

n to

ΔΩ

Ar

ΔΩAr sDIC sTA Temp Sal

−75 −70 −65 −60 −551.04

1.05

1.06

1.07

1.08

1.09

(b)

SACCF−N

PF−S

PF−M PF−N

SAF−S

SAF−M

SAF−N

TA

:DIC

Latitude [ °S]

Figure 5. From surface water measurements along a Southern

Ocean transect (a) contributions of changing sDIC (red), sTA (blue),

temperature (green), and salinity (magenta) to changing aragonite

saturation state (black, �Ar) relative to the start of the transect and

(b) TA to DIC ratios. The locations of the Subantarctic Front (SAF),

the Polar Front (PF), and the southern Antarctic Circumpolar Cur-

rent Front (SACCF) from Sokolov and Rintoul (2009) are indicated

(gray lines).

Given that SST is constant, this localized increase in �Ar is

likely due to phytoplankton photosynthesis. This region may

be along the Antarctic Slope Front that is known for higher

biological activity (Jacobs, 1991). There is another step in

�Ar from ∼ 1.4 to ∼ 1.55 between 68 and 66◦ S across the

SACCF-N. This step also corresponds with a decrease in

pCO2 from ∼ 370 to ∼ 340 µatm. Elevated POC concentra-

tions between the SACCF-N and the PF-M correspond with

these lower pCO2 values and again indicate enhanced phyto-

plankton photosynthesis. Rubin (2003) also found that pCO2

is reduced south of the PF (170◦W) due to primary produc-

tion.

To further gain insight into why �Ar increases along our

transect, we quantify the contribution of changing sDIC (cal-

culated from TA and pCO2), sTA, SST, and salinity to

changing �Ar (Fig. 5a). The dominant control is declining

sDICcalc from ∼ 2240 to ∼ 2140 µmol kg−1 between 75 and

55◦ S, which causes �Ar to increase by 0.87 if sTA, SST,

and salinity are held constant (Fig. 6). Declining sTA from

∼ 2340 to ∼ 2310 µmol kg−1 partially counters the influence

of sDICcalc and reduces �Ar by 0.28. The influences of SST

and salinity on �Ar are minimal.

�Ar variability is driven almost entirely by changes in

sDICcalc from 75◦ S to the PF-S. Between the PF-S and

the SAF-N, variability in �Ar is influenced by the oppos-

ing effects of sDICcalc and sTA. The TA :DICcalc ratio and

�Ar are constant between the PF-S and the SAF-S since

both sDICcalc and sTA decrease at the same rate (Fig. 5b).

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H. B. DeJong et al.: Carbonate saturation state in the Ross Sea and Southern Ocean 6887

2120

2140

2160

2180

2200

2220

2240

2260

sDIC

calc

[μm

ol k

g−1 ]

SACCF−NPF−S

PF−M PF−NSAF−S

SAF−MSAF−N

(a)

−75 −70 −65 −60 −552300

2310

2320

2330

2340

2350

2360

2370

sTA

[μm

ol k

g−1 ]

Latitude [ °S]

(b)

Figure 6. Measured surface water salinity normalized (a) DIC cal-

culated from pCO2, TA, temperature, and salinity and (b) TA. The

locations of the Subantarctic Front (SAF), the Polar Front (PF), and

the southern Antarctic Circumpolar Current Front (SACCF) from

Sokolov and Rintoul (2009) are indicated (gray lines).

Between the SAF-S and the SAF-N, �Ar increases since

sDICcalc declines faster than sTA . North of the SAF-N, �Ar

variability is again driven by sDICcalc. �Ar increases due to

a decrease in sDICcalc while sTA remains constant.

We examine possible controls on sDICcalc along the tran-

sect. The concentration of sDICcalc is highest south of the

PF-S due to upwelling of CDW (Fig. 6a). To evaluate

the properties of CDW, we use data from the 2011 Re-

peat Hydrography Cruise SO4P, which is part of the US

Climate Variability and Predictability (CLIVAR) program

(Swift and Orsi, 2012) (available at: http://www.clivar.org/

resources/data/hydrographic). We only use data from hydro-

casts located between 168◦ E and 73◦W where the bottom

depth is > 1000 m (Fig. 2b). We reject the data from hy-

drocast 46(B) where the deep DIC data below 200 m are

∼ 30 µmol kg−1 higher than the rest of the stations. Follow-

ing Sweeney (2003), CDW is defined as centered on the level

of maximum temperature below 150 m.

From this CLIVAR data set, CDW has a sDIC value of

2243± 3 µmol kg−1. Between 75 and 74◦ S, sDICcalc con-

centration of surface water is also 2243± 5 µmol kg−1, in-

dicating little modification to CDW and consistent with

the observation that this region was covered by sea ice

even during the summer of 2013. At 74◦ S sDICcalc drops

to ∼ 2220 µmol kg−1 and by 66◦ S, across the SACCF-N,

sDICcalc drops to ∼ 2200 µmol kg−1. This 40 µmol kg−1 de-

crease in sDICcalc between Antarctica and the PF-S is con-

sistent with the observed drops in pCO2 that we attributed

to photosynthesis. Rubin et al. (1998) also observed a 30–

50 µmol kg−1 decrease in sDIC at 67◦ S in Pacific Antarctic

waters between winter and summer that they attribute to pri-

mary productivity.

sDICcalc continues to drop from ∼ 2220 µmol kg−1 at the

PF-S to ∼ 2140 µmol kg−1 at 55◦ S, consistent with surface

DIC measurements between 70 and 40◦ S compiled by Mc-

Neil et al. (2007). There are multiple factors likely respon-

sible for this decrease in sDICcalc. Both satellite (Arrigo

et al., 2008) and in situ measurements (Reuer et al., 2007)

show that annual primary productivity increases from south

to north in the Southern Ocean. In addition, surface waters

north of the PF advect northwards and accumulate a sDIC

deficit. Finally, warmer water holds less DIC while in equi-

librium with the atmosphere. There is little net air–sea CO2

flux between 75 and 55◦ S (except for net efflux at 60◦ S)

since warming and increased biological production compen-

sate each other (Takahashi et al., 2012).

We also examine possible controls on sTA concentrations

along the transect. The concentration of sTA is also highest

south of the PF-S due to upwelling of CDW. Based off the

CLIVAR data set, the sTA of CDW is 2334± 3 µmol kg−1.

The sTA of surface water between 74◦ S and the PF-S

is ∼ 2340 µmol kg−1, slightly higher than its CDW source

(Fig. 6b). Nitrate drawdown during photosynthesis may ex-

plain the elevated sTA. Between 75 and 74◦ S, sTA exceeds

2360 µmol kg−1. One possible explanation is that ikaite

(CaCO3 · 6H2O), a mineral that has been observed directly

and indirectly to precipitate in Antarctic sea ice (Dieck-

mann et al., 2008; Fransson et al., 2011), dissolved into sur-

face waters during the summer causing sTA concentrations

to increase. Between the PF-S and SAF-N, sTA drops to

2310 µmol kg−1 where the concentrations level off. This drop

appears to be in part due to the mixing of two end member

water masses, AASW south of the PF-S and subantarctic sur-

face water north of the SAF-N. The decreasing sTA is consis-

tent with the suggestion of Millero et al. (1998) that a nega-

tive linear relationship between sTA and SST is due to colder

water being indicative of greater upwelling of TA rich water.

This data set supports the argument that increased up-

welling of CDW from strengthening westerly winds will

increase OA in the Southern Ocean (Lenton et al., 2009).

While the TA :DIC ratio for CDW is 1.040± 0.002, the

TA :DICcalc ratio for surface waters between 75◦ S and the

PF-S ranges from 1.046 to 1.064 (Fig. 5b). Therefore in-

creased upwelling will lower the TA :DIC ratio and cause

�Ar to decrease.

4.3 Estimate of wintertime surface �Ar values in the

Ross Sea

Efforts to predict winter �Ar undersaturation in the Ross Sea

are complicated by the complete lack of carbon system mea-

surements from the winter months in the Ross Sea.

McNeil et al. (2010) estimated winter surface water �Ar

by using the lowest observed �Ar value from early spring

when the Ross Sea is still covered by sea ice. They used mid-

www.biogeosciences.net/12/6881/2015/ Biogeosciences, 12, 6881–6896, 2015

6888 H. B. DeJong et al.: Carbonate saturation state in the Ross Sea and Southern Ocean

October and early November carbon system measurements

from the Joint Global Ocean Flux Study (JGOFS) (Sweeney

et al., 2000b). Although sea-ice algae productivity peaks in

November, its impact on water column DIC concentrations

is likely to be negligible (Saenz and Arrigo, 2014). McNeil

et al. (2010) found that early spring surface water �Ar was

∼ 1.2. There was a single �Ar value < 1.1 that they used as

an initial condition along with the IPCC US92a scenario to

predict that surface waters of the Ross Sea could begin to ex-

perience seasonally undersaturated conditions with respect to

aragonite as early as 2015 if full equilibrium with rising at-

mospheric CO2 is achieved. Based on a three-dimensional

coupled ice, atmosphere, and ocean model (Arrigo et al.,

2003, Tagliabue and Arrigo, 2005), McNeil et al. (2010) ar-

gued that only 35 % of the atmospheric CO2 signal equili-

brates with Ross Sea surface waters due to sea ice, upwelling

of CDW, and short residence times, thereby delaying the on-

set of aragonite undersaturation until 2045. Decadal winter-

time surface carbon system measurements do not exist to di-

rectly validate this disequilibrium assumption. In addition,

McNeil et al. (2010) would inaccurately predict when the

Ross Sea would become undersaturated with respect to arag-

onite if the minimum wintertime surface �Ar value used was

low due to measurement error.

To independently calculate �Ar from early spring sur-

face waters, we use the LDEO pCO2 measurements from

November 1994, 1997, 2005, and 2006 that are from the

Ross Shelf (defined by the 1000 m isopleth) and are south of

74◦ S (Fig. 7a). The earliest pCO2 measurements are from

16 November 1994, 17 November 1997, 6 November 2005,

and 13 November 2006 when much of the Ross Sea is still

covered in sea ice. The earliest measurements from 2005/06

are more likely to represent winter conditions since they are

from 74◦ S as the NBP entered the Ross Sea. Conversely,

the earliest measurements from 1994/97 are from the 76.5◦ S

line, close to where the Ross Sea polynya opens up from.

We calculate wintertime TA in the Ross Sea by estab-

lishing a salinity–TA relationship using data from Bates et

al. (1998), Sweeney et al. (2000b), and our own hydrocast

TA measurements from the upper 10 m (Fig. A1 in the Ap-

pendix). Since one unit of nitrate drawdown increases TA by

one unit, the TA measurements are adjusted to winter nitrate

concentrations of 29 µmol kg−1 (the mean nitrate concentra-

tion between 200 and 400 m from our cruise). The relation-

ship between TA and salinity is consistent among these inde-

pendent data sets and the standard deviation of the residuals

for TA is ±5 µmol kg−1.

We calculate historical �Ar using historical pCO2 mea-

surements, salinity-derived TA, SST, and salinity. Phos-

phate and silicate are set to the winter values of 2.1 and

79 µmol kg−1, respectively. The TSG salinity data from the

historical pCO2 measurements appear reasonable and are un-

calibrated. While the largest offset in TSG salinity compared

with Autosal measurements is 0.3, such error is not typical.

To test the possible impact of a poor salinity calibration, we

165o E 180oW 165 o

W 78

o S

76o S

74oS

72oS

(a)

5 10 15 20 25 301

1.5

2

2.5

November

ΩA

rago

nite

(b)1994199720052006

Figure 7. Estimating winter surface aragonite saturation states

(�Ar): (a) map of surface pCO2 measurements from the LDEO

pCO2 database (http://www.ldeo.columbia.edu/res/pi/CO2) used in

this study from November 1994 (blue), 1997 (red), 2005 (green),

and 2006 (black). Blue line is the 1000 m isobath. (b) aragonite

saturation state (�Ar) of surface waters from November calculated

from pCO2, salinity-derived TA, temperature, and salinity.

recalculate �Ar for all pCO2 measurements after increasing

salinity by 0.3. TA calculated from the observed TA–salinity

relationship increases by ∼ 21 µmol kg−1 and �Ar increases

by 0.024± 0.003.

The lowest �Ar measurements are 1.24 in 1994, 1.25 in

1997, 1.22 in 2005, and 1.20 in 2006 (Fig. 7b). Although

�Ar declines from 1994 to 2006, we have low confidence

in any trend due to spatial–temporal sampling biases. The

lowest �Ar values are consistently between 1.2 and 1.3 as the

ship crossed sea-ice-covered regions and open water that had

experienced DIC drawdown. With the exception of a single

measurement, the lowest 1996/97 �Ar values from McNeil

et al. (2010) are also ∼ 1.2. The similarity between the �Ar

values reported by McNeil et al. (2010) from 1996/97 and our

2005/06 values is consistent with their delayed acidification

hypothesis.

A simple calculation also suggests that wintertime �Ar

values may be closer to 1.2 than 1.1. If salinity is 34.5,

approximately the mean salinity of the water column, TA

would be 2339 µmol kg−1 based on the observed TA–salinity

linear relationship. Sweeney (2003) estimates winter pCO2

values of ∼ 425 µatm based on deep pCO2 measurements

made during early spring. Setting salinity to 34.5, TA to

2339 µmol kg−1, pCO2 to 425 µatm, temperature to −1.89,

silicate to 79 µmol kg−1, and phosphate to 2.1 µmol kg−1

yields a �Ar value of 1.22.

Although pCO2 measurements of surface waters colder

than −1.75 ◦C south of 60◦ S typically reach ∼ 410 µatm by

September, Takahashi et al. (2009) present a few measure-

ments of ∼ 450 µatm. Even if pCO2 reaches 450 µatm dur-

ing winter in the Ross Sea, �Ar would be 1.16 (with salinity

at 34.5 and TA at 2339 µmol kg−1). In order to obtain �Ar of

1.1, pCO2 would need to be∼ 480 µatm, a value that appears

unreasonably high given the available data sets from the Ross

Sea.

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H. B. DeJong et al.: Carbonate saturation state in the Ross Sea and Southern Ocean 6889

Table 2. Water properties of CDW from McNeil et al. (2010) and CLIVAR.

Data source Salinity DIC (µmol kg−1) TA (µmol kg−1) PO4 (µmol kg−1) SiO4 (µmol kg−1) �Ar

McNeil et al. (2010) 34.70± 0.02 2255± 1 2330 2.22± 0.01 93.5± 1.2 1.01

CLIVAR 34.71± 0.02 2257± 3 2348± 4 2.21± 0.04 95.6± 6.0 1.18

McNeil et al. (2010) calculated the �Ar of water arriv-

ing onto the Ross Shelf following the recipes of Jacobs et

al. (1985): 50 % CDW, 25 % Tmin water (minimum temper-

ature in upper 100 m), and 25 % AASW. To calculate the �Ar

of these three source water masses, they used hydrocast tem-

perature, salinity, and DIC data collected during the austral

winter of 1994 from north of the Ross Shelf as described in

Sweeney (2003). They calculated that the average �Ar of in-

coming water would be 1.08.

We independently calculate �Ar of incoming water us-

ing the 2011 CLIVAR hydrocast data from north of the

Ross Shelf between 168◦ E and 73◦W as described earlier

(Fig. 2b). The �Ar of water in the upper 100 m (AASW and

Tmin) from the CLIVAR data set is 1.36± 0.13 and the �Ar

of CDW (maximum temperature below 150 m) is 1.18± 0.03

(Fig. A2 in the Appendix). Even if 100 % of the incoming

water onto the Ross Shelf is CDW, the �Ar of this incom-

ing water would be greater than 1.08. While most properties

of CDW are similar between the 2011 CLIVAR data and the

1994 data used by McNeil et al. (2010), the TA of CDW from

the CLIVAR data set is 18 µmol kg−1 higher (Table 2).

Another approach to estimate the �Ar of winter surface

waters is to use the properties of water below 200 m. For the

TRACERS data, sTA below 200 m is 2338± 2 µmol kg−1.

For the JGOFS autumn cruise (NBP 97-3) sTA below 200 m

is 2339± 2 µmol kg−1. Using the CLIVAR data set, sTA of

CDW from off the Ross Shelf is 2334± 3 µmol kg−1. This

consistency between independent data sets suggests that we

can accurately estimate winter TA in the Ross Sea.

The range in sDIC below 200m is much greater than that

for sTA (Table 2). The lowest value is 2220± 5 µmol kg−1

from our cruise and the highest is 2237± 3 µmol kg−1 from

the summer JGOFS cruise (NBP 97-01). This range in sDIC

concentrations below 200 m is not surprising given that sDIC

concentrations vary across the input water masses. In addi-

tion, sDIC concentrations below 200 m will be influenced by

carbon export particularly in summer and early autumn and

over multiple seasons’ air-to-sea flux of CO2.

Assuming that deep water concentrations of TA and

DIC are relatively unmodified following wintertime deep

convective mixing, we estimate the �Ar of winter sur-

face water by setting TA to 2338 µmol kg−1, salinity to

34.5, temperature to −1.89 ◦C, phosphate to 2.1 µmol kg−1,

and silicate to 79 µmol kg−1. If DIC concentrations are

2220 µmol kg−1, �Ar would be 1.37. If sDIC concentrations

are 2237 µmol kg−1, �Ar would be 1.24 and pCO2 would be

417 µatm.

These results are consistent with a study by Matson et

al. (2014) where early spring �Ar at 20 m depth calculated

using pH and salinity-derived TA was 1.2–1.3 from Hut

Point (bottom depth > 200 m) and Cape Evans (bottom depth

< 30 m) in McMurdo Sound. In Prydz Bay, the lowest mea-

sured winter surface �Ar values were also ∼ 1.2 for both

1993–1995 (Gibson and Trull, 1999; McNeil et al., 2011) and

2010–2011 (Roden et al., 2013). Weeber et al. (2015) using

hydrocast data estimated that the �Ar of winter water in the

Weddell Sea was∼ 1.3. In the Mertz Polynya, the lowest �Ar

value at 100 m (below the mixed layer) was 1.2 (Shadwick et

al., 2013). In Arthur Harbor on the western Antarctic Penin-

sula the lowest winter surface �Ar value was 1.31 (Schram

et al., 2015).

A few studies find Antarctic winter �Ar values for surface

water below 1.2. Hauri et al. (2015) used LDEO pCO2 mea-

surements and predicted TA from salinity to estimate winter

�Ar values of surface water in the western Antarctic Penin-

sula. They found that 20 % of �Ar values were below 1.2

during the spring and winter, with a few winter values near

undersaturation. It is not surprising that winter surface �Ar

values are lower in the Antarctic Peninsula than the Ross Sea

given less sea ice in the Peninsula. In another study, Kapsen-

berg et al. (2015) report �Ar at 18 m depth (bottom depth

< 30 m) at two coastal sites in McMurdo Sound, the Jetty,

and Cape Evans, for December–May and November–June,

respectively, using pH and salinity-derived TA as input vari-

ables. The lowest �Ar observations were from May at both

sites and were 1.22 and 0.96 at the Jetty and Cape Evans. The

maximum calculated pCO2 was 559 at Cape Evans. The low

�Ar and high calculated pCO2 values measured by Kapsen-

berg et al. (2015) may represent differences between coastal

and open ocean systems – there may be a coastal amplifica-

tion signal when sinking organic matter hits a shallow bed.

Another possibility is that their carbon system time series,

particularly at Cape Evans, is inaccurate. After conditioning

and calibrating their pH measurements using discrete wa-

ter samples, for logistical reasons Kapsenberg et al. (2015)

could not collect additional validation samples during de-

ployment or measure multiple carbon system parameters for

crosscheck. Although the SeaFET pH sensors that they used

are generally stable, they can drift (Bresnahan et al., 2014).

Kapsenberg et al. (2015) have no means to assess possible

pH sensor drift.

Following McNeil et al. (2010) and a Representative Con-

centration Pathway (RCP8.5) scenario (Meinshausen et al.,

2011), we use the lowest �Ar values from 2006 (�Ar = 1.20,

www.biogeosciences.net/12/6881/2015/ Biogeosciences, 12, 6881–6896, 2015

6890 H. B. DeJong et al.: Carbonate saturation state in the Ross Sea and Southern Ocean

pCO2 = 428 µatm, TA= 2328 µmol kg−1, salinity= 34.33,

SST=−1.87 ◦C, phosphate= 2.1 µmol kg−1, sili-

cate= 79 µmol kg−1) to assess when the Ross Sea could

become corrosive to aragonite. While shelf water salinity in

the Ross Sea has declined by 0.03 decade−1 from 1958 to

2008 (Jacobs and Giulivi, 2010), we show that such rates

of change will have inconsequential effects on �Ar. For

equilibrium conditions, surface waters in the Ross Sea would

become corrosive to aragonite by 2040 (2092 for calcite)

when atmospheric CO2 concentrations exceed 485 ppm. In

the disequilibrium scenario (McNeil et al., 2010), surface

aragonite undersaturation state would occur by 2071 (2185

for calcite) when atmospheric CO2 concentrations exceed

677 ppm.

Mattsdotter Björk et al. (2014) also predicted the onset of

summertime aragonite in the Ross Sea. Their lowest �Ar

value was also ∼ 1.2 and they estimated onset of under-

saturation between 2026 and 2030 by increasing DIC by

10 µmol kg−1 per decade. This approach does not take into

account air–sea CO2 disequilibrium. In contrast, Hauck et

al. (2010) found that only 3–5 µmol kg−1 of anthropogenic

carbon accumulated per decade between 1992 and 2008 in

shelf water of the Weddell Sea. In short, our analysis suggests

that it may be possible to prevent future winter aragonite un-

dersaturation of surface waters in the Ross Sea. For instance,

CO2 concentrations never exceed 543 ppm in the CO2 stabi-

lization scenario RCP4.5 (Meinshausen et al., 2011).

If the Ross Sea experiences aragonite undersaturation dur-

ing winter in the future, live pteropod shells would start dis-

solving, making them more vulnerable to predation and bac-

terial infection (Bednaršek et al., 2012, 2014). In particular,

pteropod larvae develop during the winter/spring (Gannefors

et al., 2005; Hunt et al., 2008) and their shells have been

shown to completely dissolve within weeks of exposure to

aragonite undersaturation (Comeau et al., 2010). Declines in

pteropod populations may reduce carbon export (Manno et

al., 2010) and could have dramatic ecological effects up the

food web.

Antarctic deep sea hydrocorals may also decline or disap-

pear at the onset of aragonite undersaturation (Shadwick et

al., 2014). In addition, the shells of post-mortem bivalves and

brachiopods show significant dissolution within 2 months of

exposure to undersaturated conditions, although live organ-

isms may be able to compensate for this dissolution (McClin-

tock et al., 2009). For instance, Cummings et al. (2011) show

that the Antarctic bivalve Laternula elliptica can increase

calcification in undersaturated conditions. However, the as-

sociated energy costs may be difficult to maintain over the

long term, especially for larvae. Stumpp et al. (2012) shows

that while echinoid larvae can maintain calcification in high

pCO2 treatments, increased energetic costs reduce growth

rates and ultimately increase mortality. Larvae of the Antarc-

tic sea urchin Sterechinus neumayeri and sea star Odontaster

validus are smaller and exhibit abnormal development un-

der elevated pCO2 treatments (Byrne et al., 2013; Gonzalez-

Bernat et al., 2013; Yu et al., 2013). In addition, the syner-

gistic effects of warming and OA could impact echinoderm

fertilization and embryo development (Ericson et al., 2012).

Although it is not clear to what extent species may acclima-

tize or adapt (e.g., Suckling et al., 2015), the onset of arag-

onite undersaturation during winter months may have pro-

found impacts on the Ross Sea ecosystem.

5 Conclusions

Our study demonstrates the possibility of setting up under-

way TA measurement systems. Although our system was rel-

atively unattended, carbon system crosschecks and compar-

isons between hydrocast and underway data indicate that our

measurements were accurate. Similar underway TA systems

could be set up on scientific vessels and ships of opportunity

in undersampled regions of the world’s oceans.

We find that the seasonal increase in �Ar in the Ross Sea

by early autumn is driven almost entirely by phytoplankton

photosynthesis. In the Southern Ocean between the Ross Sea

and Chile we find that �Ar also increases mainly due to de-

clining DICcalc although declining TA partially counters the

influence of declining DICcalc. The influences of SST and

salinity on �Ar are minimal in the Ross Sea and on our

Southern Ocean transect.

We establish a salinity–TA relationship for the winter that

is consistent across independent data sets. Using historical

pCO2 measurements from early spring along with TA pre-

dicted from salinity, we argue that it is unlikely that the Ross

Sea actually experienced winter surface �Ar values of ∼ 1.1

during 1996 (as per McNeil et al., 2010) and that a �Arvalue

of ∼ 1.2 may more accurately represent current winter con-

ditions.

Since predictions are sensitive to current surface winter-

time �Ar values as well as the extent of disequilibrium,

highly accurate carbon system measurements from the winter

are crucial. It is also essential to measure more than two car-

bon system parameters for crosscheck. For instance, pH and

pCO2 sensors on moorings and floats could be used with TA

predicted from salinity to calculate � during the winter.

Our analysis indicates that the Ross Sea will not experi-

ence aragonite undersaturation until the year 2070 follow-

ing RCP8.5. In some CO2 stabilization scenarios, including

RCP4.5 (Meinshausen et al., 2011), the Ross Sea may avoid

becoming corrosive to aragonite.

Biogeosciences, 12, 6881–6896, 2015 www.biogeosciences.net/12/6881/2015/

H. B. DeJong et al.: Carbonate saturation state in the Ross Sea and Southern Ocean 6891

Appendix A

33.5 34 34.5 352240

2260

2280

2300

2320

2340

2360

TA

(μm

ol k

g−1 )

Salinity

TA = 68.9*Sal − 37.8

R2 = 0.94

Figure A1. Linear regression between TA and salinity with surface

data from February to March 2013 (blue; this study), November

to December 1994 (green; Bates et al., 1998), December to Jan-

uary 1995/1996 (red; Bates et al., 1998), and April 1997 (magenta;

Sweeney et al., 2000a). TA has been corrected to a nitrate concentra-

tion of 29 µmol kg−1 to account for the effects of nitrate drawdown

on TA (Brewer and Goldman, 1976).

0.8 1 1.2 1.4 1.6 1.8

0

200

400

600

800

1000

Dep

th (

m)

ΩAragonite

ΩAr

= 1.08 →

Figure A2. Profiles of aragonite saturation state (�Ar) from off the

Ross Shelf (see Fig. 2b) from the CLIVAR program (NBP 11-02)

calculated from TA, DIC, temperature, and salinity at surface pres-

sures.

www.biogeosciences.net/12/6881/2015/ Biogeosciences, 12, 6881–6896, 2015

6892 H. B. DeJong et al.: Carbonate saturation state in the Ross Sea and Southern Ocean

The Supplement related to this article is available online

at doi:10.5194/bg-12-6881-2015-supplement.

Acknowledgements. This work was supported by the US NSF

(OPP-1142044 to R. B. Dunbar) and a NSF graduate research

fellowship grant (DGE-114747 to H. B. DeJong). We thank the

captain and crew of the R/V Nathaniel B. Palmer. We are grateful to

S. Bercovici for the nitrate data. The comments of two anonymous

reviewers greatly improved this paper.

Edited by: L. Bopp

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